PALEOCEANOGRAPHY, VOL. 28, 91–104, doi:10.1002/palo.20014, 2013
Deglacial development of (sub) sea surface temperature and
salinity in the subarctic northwest Pacific: Implications for
upper-ocean stratification
Jan-Rainer Riethdorf,1,3 Lars Max,2 Dirk Nürnberg,1 Lester Lembke-Jene,2
and Ralf Tiedemann2
Received 25 June 2012; revised 13 January 2013; accepted 17 January 2013; published 23 March 2013.
[1] Based on models and proxy data, it has been proposed that salinity-driven stratification
weakened in the subarctic North Pacific during the last deglaciation, which potentially
contributed to the deglacial rise in atmospheric carbon dioxide. We present high-resolution
subsurface temperature (TMg/Ca) and subsurface salinity-approximating (d18Oivc-sw) records
across the last 20,000 years from the subarctic North Pacific and its marginal seas, derived
from combined stable oxygen isotopes and Mg/Ca ratios of the planktonic foraminiferal
species Neogloboquadrina pachyderma (sin.). Our results indicate regionally differing
changes of subsurface conditions. During the Heinrich Stadial 1 and the Younger Dryas
cold phases, our sites were subject to reduced thermal stratification, brine rejection due to
sea-ice formation, and increased advection of low-salinity water from the Alaskan Stream.
In contrast, the Bølling-Allerød warm phase was characterized by strengthened thermal
stratification, stronger sea-ice melting, and influence of surface waters that were less diluted
by the Alaskan Stream. From direct comparison with alkenone-based sea surface
temperature estimates (SSTUk0 37), we suggest deglacial thermocline changes that were
closely related to changes in seasonal contrasts and stratification of the mixed layer. The
modern upper-ocean conditions seem to have developed only since the early Holocene.
Citation: Riethdorf, J.-R., L. Max, D. Nürnberg, L. Lembke-Jene, and R. Tiedemann (2013), Deglacial development of
(sub) sea surface temperature and salinity in the subarctic northwest Pacific: Implications for upper-ocean stratification,
Paleoceanography, 28, 91–104, doi:10.1002/palo.20014.
salinity-driven stratification, the exchange of gas, heat, and
nutrients between deep and surface water is limited in the N
Pacific. In contrast, the modern Southern Ocean releases
carbon dioxide (CO2) to the atmosphere due to upwelling of
deep waters. This led to the assumption that high-latitude
ocean stratification drives changes in atmospheric CO2
concentrations during recent glacial-interglacial cycles [Haug
et al., 1999; Sigman and Boyle, 2000; Sigman and Haug,
2003; Sigman et al., 2004, 2010; Jaccard et al., 2005]. However, growing paleoceanographic evidence [e.g., Okazaki et
al., 2010] indicates that during the last deglaciation, deep water was formed in the N Pacific and that the regional halocline
was not a permanent feature. Hence, the N Pacific might have
contributed to the deglacial rise of atmospheric CO2.
[3] High-resolution records depicting the deglacial paleoceanographic evolution in the subarctic N Pacific are sparse
due to a shallow lysocline and corrosive bottom waters
limiting CaCO3 preservation. Available reconstructions of
sea surface temperature (SST) and salinity indicate strong
oceanographic changes and climate oscillations in the
subarctic N Pacific realm during the last deglaciation
(20–10 ka BP) [e.g., Ternois et al., 2000; Sarnthein et al.,
2004, 2006; Gebhardt et al., 2008; Sagawa and Ikehara,
2008; Caissie et al., 2010], which are similar to those recorded
in Greenland ice [Grootes et al., 1993; NGRIP members,
2004]. These are the cold periods of the Heinrich Stadial 1
1. Introduction
[2] No deep water is formed in the modern subarctic
North Pacific (N Pacific). There, a relatively steep salinity
gradient (permanent halocline) prevents surface water from
sinking, isolating it from the underlying nutrient-rich deep
water [Haug et al., 1999]. In the N Pacific, upper-ocean
stratification most likely developed around 2.7 million years
ago [Haug et al., 1999, 2005; Sigman et al., 2004] and is
maintained by a restricted meridional exchange between
subpolar and subtropical waters, atmospheric low-latitude
moisture transport from the Atlantic to the Pacific, and
northward moisture flux by the Asian monsoon [Warren,
1983; Emile-Geay et al., 2003]. As a consequence of
Additional supporting information may be found in the online version of
this article.
1
Helmholtz Centre for Ocean Research Kiel (GEOMAR), Kiel, Germany.
2
Alfred Wegener Institute for Polar and Marine Research, Bremerhaven,
Germany.
3
Department of Ocean Floor Geoscience, Atmosphere and Ocean
Research Institute, University of Tokyo, Kashiwa, Chiba, Japan.
Corresponding author: Jan-Rainer Riethdorf, Helmholtz Centre for
Ocean Research Kiel (GEOMAR), Wischhofstr. 1-3, D-24148 Kiel, Germany. ([email protected])
©2013. American Geophysical Union. All Rights Reserved.
0883-8305/13/10.1002/palo.20014
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
stratification in the N Pacific is a relatively recent feature
as suggested by Sarnthein et al. [2004].
(H1, 18.0–14.7 ka BP) [Sarnthein et al., 2001] and the
Younger Dryas (YD, ~12.9–11.7 ka BP) [Blockley et al.,
2012], and the warm phases of the Bølling-Allerød (B/A,
~14.7–12.9 ka BP) [Blockley et al., 2012] and the Preboreal
(PB, ~11.7–11.0 ka BP). Recent studies also found evidence
for deep water ventilation in the NW Pacific during H1 and
the YD [Ahagon et al., 2003; Ohkushi et al., 2004; Sagawa
and Ikehara, 2008; Okazaki et al., 2010], while at the same
time the Atlantic meridional overturning circulation (AMOC)
collapsed or declined [McManus et al., 2004]. In contrast, during the B/A, higher ventilation ages have been found in the
western N Pacific, which is indicative of reduced deep water
ventilation [Okazaki et al., 2010]. In agreement with this observation, general circulation models predict a strengthening
of the Pacific meridional overturning circulation (PMOC),
which results from a rise in salinity in the N Pacific due to a
weakened AMOC [e.g., Menviel et al., 2012]. These studies
controversially argue either for an atmosphere-controlled
“in-phase” [Mikolajewicz et al., 1997; Krebs and Timmermann,
2007; Okumura et al., 2009] or for an ocean-controlled
“antiphase” [Schmittner et al., 2003, 2007; Saenko et al.,
2004] relationship between the thermal evolution of the North
Atlantic (N Atlantic) and the N Pacific. The antiphase relationship is supported by SST and salinity reconstructions from
planktonic foraminifera at Site MD01-2416 and ODP Site
883 indicating that during H1, maxima in SST were accompanied by increased salinity [Kiefer et al., 2001; Sarnthein et al.,
2006; Gebhardt et al., 2008]. Higher salinity during H1 is
supported by Mg/Ca- and d18O-based results from core
GH02-1030 off Japan [Sagawa and Ikehara, 2008], which
also argues for a potential disappearance of the halocline.
The results from core MD01-2416, however, are in conflict
with alkenone-based SST reconstructions from the N Pacific
realm, which are more supportive of the in-phase models.
Results indicate restricted marine productivity during H1 and
a SST maximum that occurred during the subsequent B/A in
the NE Pacific [Kienast and McKay, 2001; Barron et al.,
2003], the Bering Sea [Caissie et al., 2010; Max et al.,
2012], and the Okhotsk Sea [Ternois et al., 2000; Harada
et al., 2006a; Seki et al., 2009]. Mg/Ca-based results from core
GH02-1030 also show a rise in SST during the B/A [Sagawa
and Ikehara, 2008]. In summary, these observations may
point to a regionally differing development of the thermocline,
because sites MD01-2416 (Detroit Seamount, open subarctic
NW Pacific) and GH02-1030 (Kuroshio-influenced NW
Pacific) provide quite different oceanographic settings with
respect to the N Pacific marginal seas. Consequently,
additional reconstructions of SST and salinity, which allow
for direct comparisons between alkenone- and Mg/Ca-based
results, are essential to understand changes in the structure of
the upper water column and the SST development of the
subarctic N Pacific.
[4] Here, we report combined stable oxygen isotope and
Mg/Ca-based reconstructions of subsurface temperatures
(TMg/Ca) and d18Oivc-sw (approximating subsurface salinity)
from sediment cores recovered from the southern Okhotsk
Sea, the NW Pacific off Kamchatka, and the western Bering
Sea covering the last 20 ka BP. Our results, which are compared to alkenone-based SST estimates (SSTUk0 37) derived
from the same samples [Max et al., 2012], indicate deglacial
variations in mixed layer stratification (MLS). Moreover, we
present supporting evidence that the modern salinity-driven
2. Regional Setting
[5] The subarctic N Pacific is characterized by a largescale cyclonic surface circulation pattern (Figure 1). At
~40 N, the North Pacific Current, an extension of the
subtropical Kuroshio Current, flows eastward and brings
relatively warm water (~10 C) into the Alaskan gyre in the
NE Pacific. From here, the Alaskan Current, fed by freshwater discharge from the North American continent [Kowalik
et al., 1994; Weingartner et al., 2005], transports surface
water to the north. Subsequently, the Alaskan Stream flows
westward along the Aleutian Island Arc, thereby causing
surface water to flow into the Bering Sea through several
passes. Within the Bering Sea, a cyclonic surface circulation
develops with the East Kamchatka Current (EKC) and the
Bering Slope Current acting as boundary currents. Cold
and nutrient-rich surface waters leave the Bering Sea
through Bering Strait into the Arctic Ocean, but the main
outflow occurs back into the NW Pacific via Kamchatka
Strait [e.g., Stabeno et al., 1999]. The northern straits of
the Kurile Islands provide inflow of Pacific water from the
EKC into the Okhotsk Sea [e.g., Katsumata and Yasuda,
2010]. In the Okhotsk Sea, the surface circulation pattern
is also cyclonic. Notably, brine rejection due to sea-ice formation leads to the production of Okhotsk Sea intermediate
water (OSIW), which determines the potential density (s0)
surface of NPIW in the N Pacific. OSIW flows out of the
Okhotsk Sea through the Kurile Straits, thereby mixing with
waters from the Western Subarctic Gyre [Yasuda, 1997;
You, 2003]. The Oyashio Current transports this relatively
cold (~4 C), low-salinity (~33 psu) water along the
Kurile Islands to the east of Japan, where it meets with
warmer and saltier water (~34–35 psu) from the Kuroshio.
Cabbeling of these water masses then forms NPIW.
[6] Characteristic oceanographic features of the subarctic
NW Pacific are the strong seasonal variability of SST and
the marked upper-ocean stratification due to the buoyant
low-salinity surface layer (Figure 2). Both are linked to the
seasonal interplay between the Siberian High and the
Aleutian Low, which in the Okhotsk and Bering seas leads
to intense winter mixing and to seasonal sea-ice formation
[e.g., Niebauer et al., 1999]. Especially during winter, the
sea-ice cover is significantly expanded in both marginal seas
and in the vicinity of the eastern Kamchatka continental
margin. Sea-ice formation releases brines, which sink in
the water column resulting in increased subsurface salinity.
During summer, MLS arises from increased insolation and
melting sea-ice, while a temperature minimum layer
(dichothermal layer) remains at ca. 100 m (Figure 2). Waters
from this layer are believed to be formed during winter in the
Bering and Okhotsk seas and to be subsequently exported to
the NW Pacific [Ohtani et al., 1972; Miura et al., 2002]. At
our study sites, the modern seasonal thermo- and
pycnoclines lie within 0–100 m. The permanent halocline
lies deeper (~100–200 m), and it further deepens toward
the northern part of the Bering Sea, as do the mixed and
the dichothermal layers (Figure 2). Notably, the calcite saturation horizon (CSH) at our sites lies between 150 and
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
Figure 1. Bathymetric map with 250 m isobathe of the subarctic NW Pacific, Okhotsk Sea, Bering Sea,
and NE Pacific (inlet). Red dots indicate sediment cores studied here, blue dots denote published reference
records: MD01-2416 and ODP Site 883 [Sarnthein et al., 2004, 2006; Gebhardt et al., 2008], and PC13
[Brunelle et al., 2010] from the NW Pacific, MD02-2489 from the NE Pacific [Gebhardt et al., 2008],
GH02-1030 off Japan [Sagawa and Ikehara, 2008], GGC27 from the Okhotsk Sea [Brunelle et al.,
2010], HLY-02-02-17JPC from Bowers Ridge [Brunelle et al., 2007, 2010], and MR06-04-PC23A from
the upper slope of the NE Bering Sea [Rella et al., 2012]. CTD stations from R/V Sonne expedition
SO201-2 (white triangles) [Dullo et al., 2009], WOA 2009 stations (white squares) [Locarnini et al.,
2010], and stations from the World Ocean Circulation Experiment (WOCE; black dots) referred to in
the text are included. The white star marks the position of Okhotsk Sea station MU’6 [Yamamoto et al.,
2001]. The general surface circulation pattern [after Tomczak and Godfrey, 1994; Stabeno et al., 1999]
is indicated by red arrows, and the dotted line marks the modern average winter sea-ice extent (after Niebauer et al., 1999; IRI/LDEO Climate Data Library, http://iridl.ldeo.columbia.edu/). AC = Alaskan Current, EKC = East Kamchatka Current, ANSC = Aleutian North Slope Current. This map was generated
with “Ocean Data View” [Schlitzer, 2011].
(color b*) and X-ray fluorescence (XRF) core logging data for
intercore correlations, as well as accelerator mass spectrometry (AMS) radiocarbon dating of planktonic foraminifera.
The stratigraphic framework of all cores including the
AMS-14C dating results is presented in detail in Max et al.
[2012] and in the Supplementary Information.
300 m, which has implications for our carbonate-based
proxy reconstructions (Supplementary Information).
3. Material and Methods
3.1. Sedimentology and Age Models
[7] This study is based on piston cores SO201-2-12KL, 77KL, -85KL, and -101KL recovered during R/V Sonne
cruise SO201 KALMAR Leg 2 in the NW Pacific off Kamchatka and in the western Bering Sea [Dullo et al., 2009]
(Table 1). Sediments are dominated by siliciclastic material
of mainly clay and silt size, which are intercalated by thin
layers of diatomaceous ooze/silt. Additional samples stem
from core LV29-114-3 retrieved from the southern Okhotsk
Sea during cruise LV29 KOMEX Leg 2 with R/V Akademik
Lavrentyev [Biebow et al., 2003] (Table 1). In this core, terrigenous sediments are overlain on top by a 175 cm thick
layer of diatomaceous sediment. Sites 12KL and 114-3 are
influenced by the EKC. All cores contain low contents of
CaCO3 (<5 wt %). However, all cores show increased
CaCO3 contents of up to 30 wt % during the B/A.
[8] Age models are based on a combined chronostratigraphic approach including high-resolution spectrophotometric
3.2. Stable Isotope and Mg/Ca Analyses
[9] Combined stable oxygen isotope (d18O) and Mg/Ca
analyses were performed on ~500 mg (~100–150 tests) of
the polar to subpolar subsurface–dwelling planktonic
foraminifer Neogloboquadrina pachyderma (sin.) (referred
to as Nps hereafter), selected from the 125–250 mm size fraction. We focused on the most abundant four-chambered
specimen of Nps. Abundance of foraminiferal tests was sufficient in all sediment cores, except for core 12KL, which
lacked sufficient amounts after 6 ka BP. d18O was measured
using a MAT 253 mass spectrometer (Thermo Scientific,
Germany) coupled with a Kiel IV Carbonate device (Thermo
Scientific, Germany). Results were calibrated to the VPDB
scale and referenced to the NBS19 standard. Analytical
long-term precision (n > 1000 samples) of the used inhouse carbonate standard (Solnhofen limestone) was
93
RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
Figure 2. (A, B) Modern seasonal profiles of in situ temperature, salinity, and potential density (s0)
during boreal winter (January-March, thin lines) and summer (July-September, thick lines), as well
as actual seawater d18Osw data (C) for the southern Okhotsk Sea, the subarctic NW Pacific, and
the western Bering Sea. Stations were chosen from WOA 2009 (stations 33545, 34115, 34548)
[Locarnini et al., 2010] (cf. Figure 1). In (B), summer conditions are shown together with CTD data
of R/V Sonne expedition SO201-2 from September 2009 (stations SO201-2-2CTD, -67CTD, 110CTD) [Dullo et al., 2009]. d18Osw data from Okhotsk Sea station MU’6 are from Yamamoto et
al. [2001]. The habitat of the planktonic foraminifer N. pachyderma (sin.) is assumed to lie in 50–
100 m (gray-shaded areas) and to be associated with an isopycnal layer. The modern hydrographic
ranges of the photic zone (0–30 m) and of the subsurface in 50–100 m are represented by red and
gray bars, respectively, next to the abscissae. Maximum mixed layer depths (dashed line) are inferred
from Miura et al. [2002] and fit WOA and CTD data. Note the thicker mixed and dichothermal layer
at Bering Sea station 110CTD.
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
from the comparison of the foraminiferal d18O signal to
what would be predicted for the d18O signal of seawater
(d18Osw) when using the hydrographic temperature data
(Supplementary Information).
[12] It has been suggested that in the NW Pacific, Nps
represents late spring and late summer/early fall temperatures [Sarnthein et al., 2006; Gebhardt et al., 2008], which
is supported by sediment trap studies [Takahashi et al.,
2002; Kuroyanagi et al., 2012]. Results for the southern
Bering Sea and the NW Pacific Subarctic Front reveal that
the main peak in CaCO3 depositional flux occurs during late
summer/early fall [Takahashi et al., 2002; Mohiuddin et al.,
2005]. We therefore assume that at our study sites, which are
additionally influenced by winter sea-ice, the main Mg/Ca
signal of Nps likely represents late summer/early fall. This
is supported by our reconstructed average Holocene (<8 ka
BP) TMg/Ca estimates of 3–4 C, which is only 1–2 C warmer
than during modern summer and fall at 50–100 m. In
contrast, modern winter (January-March) and spring temperatures lie below 3 C in 0–100 m (Figure 2).
[13] Our Mg/Ca results might be partly influenced by
dissolution of the foraminiferal tests in conjunction with
selective removal of Mg2+ ions [e.g., Regenberg et al.,
2006, and references therein], especially since our sediment
cores were recovered below the modern CSH. Variations
in Mg/Ca might thus be related to changes in CaCO3 preservation. However, we refrained from correcting the initial
Mg/Ca values as clear relationships between CaCO3
contents and foraminiferal Mg/Ca were not found and correction procedures for contamination/dissolution effects were
considered unsuitable (Supplementary Information).
Table 1. Site Information
Core
LV29-114-3
SO201-2-12KL
SO201-2-77KL
SO201-2-85KL
SO201-2-101KL
Latitude
Longitude
Depth
(mbsl)
Recovery
(m)
49 22.54’N
53 59.47’N
56 19.83’N
57 30.30’N
58 52.52’N
152 53.23’E
162 22.51’E
170 41.98’E
170 24.77’E
170 41.45’E
1765
2145
2135
968
630
9.64
9.05
11.78
18.13
18.32
0.06%. Cleaning of foraminiferal tests for Mg/Ca analyses
followed the protocol of Barker et al. [2003] and included
an additional reductive cleaning step. Samples were measured
on an axial viewing ICP-OES (VARIAN 720-ES) with an analytical long-term precision of 0.1 mmol mol-1 for Mg/Ca of
the ECRM752-1 standard. For core 114-3, earlier measurements were included in our data set to improve sampling resolution. These measurements used 30 specimens of Nps
(150–250 mm size fraction) to determine d18O and 50 specimens for Mg/Ca analyses. Cleaning also followed the protocol of Barker et al. [2003], but did not include reductive
cleaning. A radially viewing ICP-OES (Ciros CCD SOP,
Spectro A.I.) was used to determine Mg/Ca. Method details
are given in the Supplementary Information.
3.3. Mg/Ca Temperature Signal of Nps
[10] As there is no locally established temperature-Mg/Ca
calibration for Nps in the subarctic N Pacific, we applied the
linear equation of Kozdon et al. [2009]. It is based on Holocene core-top samples from high-latitude Nordic Seas used
in a cross-calibration approach between Mg/Ca and independent d44/40Ca measurements.
1
Mg=Ca mmol mol
¼ 0:13Temp:ð CÞ þ 0:35
3.4. Combined Use of Mg/Ca- and Alkenone-Based
Temperatures
[14] We compared our TMg/Ca records with the SSTUk0 37
records of Max et al. [2012], which were derived from the
same samples. Max et al. [2012] used the calibration of
Müller et al. [1998], which is widely used in N Pacific
SST reconstructions providing a standard error of 1.5 C.
For the NW Pacific and both marginal seas, it has been
shown that alkenones are mainly synthesized by Emiliania
huxleyi during late summer to early fall within the upper
50 m [Harada et al., 2003, 2006b; Seki et al., 2007]. At
our sites, surface sediment SSTUk0 37 of 5–7 C (not shown)
falls within 1–2 C when compared with the modern instrumental range at 0–50 m (Figure 2).
[15] Differences between SSTUk0 37 and TMg/Ca can result
from seasonal bias [e.g., Leduc et al., 2010]. However, for
our sites we assume a restriction of both, coccolithophores
and planktonic foraminifera, to the sea-ice-free late
summer/early fall season. Nevertheless, we consider that a
seasonal bias might have influenced the proxy records under
conditions of reduced regional sea-ice influence, e.g., during
the B/A. Accordingly, a high temperature gradient (ΔT)
between SSTUk0 37 and TMg/Ca is thought to reflect enhanced
thermal MLS. However, the combined uncertainty for
SSTUk0 37 and TMg/Ca is substantial (2.3 C), and significant
temporal trends in the records are needed for interpretation
of ΔT. In addition, ΔT might be influenced by intensified
insolation as well as by variations in the exchange of heat
between the upper and the deep ocean [Kohfeld et al.,
1996; Andersson et al., 2010].
(1)
[11] Considering the slope in equation (1), the analytical
precision for Mg/Ca translates into an uncertainty of
0.8 C. Results of Kozdon et al. [2009] indicated variable
calcification depths of Nps that are associated with an
isopycnal layer. Most other temperature calibrations for
Nps are of exponential character and assume constant calcification depths. We consider the habitat of N Pacific Nps
likely to be related to the seasonal thermo- and pycnoclines,
similar to the Nordic Seas and the Arctic Ocean. Other studies conducted there showed that shell calcification of Nps
mostly occurs at or close to the thermocline in 50–200 m
[Kohfeld et al., 1996; Bauch et al., 1997; Simstich et al.,
2003]. Based on these studies, Sarnthein et al. [2004,
2006] assumed a depth range of 30–100 m in the NW
Pacific. For the Okhotsk Sea, Bauch et al. [2002] suggested
average calcification depths at 50–200 m and concluded that
Nps lives at the bottom of the thermocline. Tow samples
from the NW Pacific showed that the habitat of Nps lies below the pycnocline (>20 m) [Kuroyanagi and Kawahata,
2004]. Accordingly, we assume that the habitat of Nps lies
at 50–100 m at our sites. Considering data from the World
Ocean Atlas (WOA), this depth range lies below the summer
(July-September) pycnocline, extends to the bottom of the
summer thermocline, and is associated with an isopycnal
layer (s0 26.4 kg m-3; Figure 2). Additional support for
the assumption of calcification at subsurface levels comes
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
amplitude variations and a TMg/Ca maximum of ~9 C. Alkenone concentrations before 15 ka BP are below detection
limit and afterward the records reflect successively higher
SSTUk0 37 from the north to the south of ~3 C [Max et al.,
2012]. No meridional gradients are found for TMg/Ca. Relative temperature changes in SSTUk0 37 and TMg/Ca consistently show a warming from H1 into the B/A and a cooling
from the B/A into the YD at all sites (Figure 3A). The deglacial SSTUk0 37 evolution almost parallels the thermal evolution registered in the NGRIP ice core from Greenland,
whereas the TMg/Ca evolution does not. Most notably, TMg/
Ca and SSTUk0 37 records in part show different trends. All
cores show TMg/Ca values of ~3–4 C during the H1 cold
phase, but only at Site 12KL a significant cooling of ~3 C
is recorded with respect to the Last Glacial Maximum
(LGM) (Figure 3A). At the end of H1 within only 1–2 kyr,
TMg/Ca increase by 2–4 C to maxima of ~5–6 C at all sites,
but to ~9 C at Site 12KL. SSTUk0 37 also increase by 2–6 C
to maxima of 6–8 C [Max et al., 2012]. While Bering Sea
cores 85KL and 101KL recorded almost similar TMg/Ca and
SSTUk0 37 until the PB, cores 114-3 and 77KL show ΔT
values of ~2–3 C during the B/A (Figure 4). Notably, core
12KL features negative ΔT values that are minimal (-6 C)
at the onset of the B/A and subsequently increasing. Following the B/A, cores 114-3, 12KL, and 77KL recorded a TMg/
Ca cooling to minima of ~3–4 C recorded during the YD.
Core 85KL still exhibits a strong Mg/Ca variability, but on
average decreasing values since the B/A. SSTUk0 37 also decrease by 2–5 C into the YD [Max et al., 2012]. During
the YD, ΔT is reduced to about +2 C at sites 114-3 and
77KL, to 0 C at Site 85KL, and it again becomes negative
(-2 C) at Site 12KL (Figure 3A). At Site 12KL, a ~2 C
warming from the YD into the PB is recorded with a pronounced TMg/Ca maximum of ~6 C (~11.5 ka BP), whereas
Bering Sea core 77KL shows TMg/Ca of ~4 C, and Okhotsk
Sea core 114-3 does not show significantly increased TMg/
Ca. The PB is subject to or followed by a 1–2 C cooling to
TMg/Ca of ~3–4 C at all sites, which remain almost constant
afterward. SSTUk0 37 increase by up to 5 C subsequent to the
YD and culminate in maxima of 9–10 C between 11 and
9 ka BP [Max et al., 2012]. The early Holocene SSTUk0 37
maximum and the summer insolation maximum occur simultaneously (Figure 3A). Consequently, both temperature
proxies significantly diverge during or shortly after the PB
until ~10 ka BP, with ΔT maxima of up to 6 C. Notably,
ΔT at Site 12KL has become positive only since the PB. Holocene TMg/Ca values are lower than those of the B/A and
YD, but are comparable to H1 values. For the last 9 kyr,
records from cores 114-3, 12KL, and 77KL point to a
gradual ~2 C decrease in SSTUk0 37, which in core 114-3 is
interrupted between 9 and 7 ka BP by a cooling to YD levels
[Max et al., 2012]. Hence, the ΔT evolution during the
middle to late Holocene overall follows the SSTUk0 37 signal
and records a ΔT minimum between 9 and 7 ka BP in core
114-3 (Figure 3A).
3.5. Ice Volume–Corrected d18Osw and the Influence of
Brines
[16] We used ice volume–corrected estimates of the local
seawater oxygen isotopic composition (d18Oivc-sw; reported
in % versus SMOW) to approximate past changes in local
salinity. d18Osw was calculated by applying the relationship
of Shackleton [1974]. Finally, we corrected for the global
ice-volume signal following Waelbroeck et al. [2002]. The
conversion of d18Osw into salinity relies on regional calibrations. Such calibrations exist for the Okhotsk Sea and the
western subarctic Pacific [Yamamoto et al., 2001, 2002].
However, the uncertainty of the d18Osw approach (~0.3%)
already translates into a salinity error of ca. 0.8 when applying the calibration of Yamamoto et al. [2001]. Since our study
sites show seasonal salinity variations of ~0.4–0.7 at 0–100 m
(Figure 2), we did not apply a respective conversion. A shift in
d18Oivc-sw toward more positive values, hence, is thought to
reflect a rise in local salinity, and vice versa.
[17] In the high latitudes, the salinity-d18Osw relationship is
not linear due to sea-ice-related brine rejection [e.g., Bauch
et al., 1995, 1997; Yamamoto et al., 2002; Hillaire-Marcel
et al., 2004]. In general, sea-ice formation in the surface layer
produces salty but isotopically light (18O-depleted) brines,
which sink in the water column and mix with subsurface
waters. Conversely, the surface layer is provided with
low-salinity, 18O-enriched water as a result of melting seaice [e.g., Hillaire-Marcel and de Vernal, 2008]. Isotopic
excursions recorded in Nps from Arctic Ocean sediments
were used to infer variations in sea-ice growth during
Heinrich events [Hillaire-Marcel and de Vernal, 2008].
For the Okhotsk Sea, a mean d18Osw deviation of
0.36 0.02% to what would be expected from the salinity-d18Osw relationship of Yamamoto et al. [2001] was
found and ascribed to the influence of brines [Yamamoto
et al., 2002]. Hence, our d18Oivc-sw records might not unequivocally reflect past salinity changes. During winter,
the northern part of the western Bering Sea is covered with
sea-ice [Niebauer et al., 1999; Zhang et al., 2010], which
might result in a stronger impact of brine rejection/sea-ice
melt on the d18Oivc-sw records toward our northern Bering
Sea sites.
3.6. Biogenic Opal and CaCO3
[18] Biogenic opal was measured on freeze-dried bulk
sediment samples (20 mg) via molybdate-blue spectrophotometry [Müller and Schneider, 1993], and concentrations
were calculated after DeMaster [1981] with a reproducibility
of 1–2 wt %. Total carbon (TC) and total organic carbon
(TOC) concentrations were determined using a Carlo Erba
CNS Analyzer (model NA-1500), with TOC being measured on previously decalcified samples. Concentrations of
CaCO3 were indirectly calculated as the difference between
TC and TOC, multiplied by 8.333.
4. Results
4.2. Reconstruction of d18Oivc-sw
[20] Reconstructed d18Oivc-sw values range between -1%
and +1% during the last 20 kyr, with the most prominent
variability also being recorded at Site 12KL (Figure 3B).
Notably, Bering Sea core 85KL and Okhotsk Sea core
4.1. Temperature Reconstructions
[19] Figure 3A shows our TMg/Ca results for the last 20 kyr
together with the SSTUk0 37 records from Max et al. [2012].
Reconstructed TMg/Ca have a similar range in all cores of
about 2 to 6 C. Core 12KL recorded the most pronounced
96
RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
Figure 3. Temperature reconstructions (A) and ice volume–corrected seawater d18O estimates (d18Oivc-sw; B)
over the last 20 kyr from the Bering Sea (SO201-2-101KL, -85KL, -77KL), the subarctic NW Pacific off
Kamchatka (SO201-2-12KL), and the southern Okhotsk Sea (LV29-114-3). The NGRIP ice core oxygen
isotope record (NGRIP members, 2004; GICC05 time scale, Rasmussen et al., 2006) is for reference. Running
five-point-averages (thick black lines) are given to smooth the records of TMg/Ca and d18Oivc-sw, and the
respective uncertainties are indicated. Relative sea level (RSL, green line) is from Waelbroeck et al. [2002]
and the green vertical bar marks the interval when RSL reached the approximate sill depth of the Bering Strait
(~50 m). Pale orange and pale blue shadings represent the B/A and PB, and H1 and YD, respectively. In (A), the
TMg/Ca records (light gray lines) are shown together with the alkenone-based SSTUk0 37 records (thin red lines)
from Max et al. [2012], plotted on the same scale. Average middle to late Holocene (<8 ka BP) temperature
estimates of 3–4 C are highlighted (gray vertical bars), while the modern photic zone (0–30 m) and subsurface
(50–100 m) temperature ranges are represented by red and gray vertical bars, respectively, next to the temperature axes (based on WOA data) [Locarnini et al., 2010] (cf. Figure 2). Mean insolation for boreal summer
(July-September) at 65 N (thick gray line) was calculated after Laskar et al. [2004]. In (B), trends toward
heavier (lighter) d18Oivc-sw signatures are equivalent to increasing (decreasing) local subsurface salinity. Toward
the northern Bering Sea sites, the d18Oivc-sw signal is assumed to be additionally influenced by 18O-enriched
meltwater during sea-ice melt and 18O-depleted brines during sea-ice formation. Dashed lines indicate
d18Oivc-sw = 0% SMOW. Red and gray vertical bars next to the d18Oivc-sw axes mark the modern d18Osw ranges
for the photic zone (0–30 m) and subsurface (50–100 m), respectively, at Okhotsk Sea station MU’6 [Yamamoto
et al., 2001]. Notably, all records are characterized by negative average d18Oivc-sw values after the PB.
d18Oivc-sw is negative at sites 114-3 and 101KL (-0.3% to
0%). During H1, Site 12KL is marked by a d18Oivc-sw
minimum of about -0.4% (Figure 3B). A similar significant
extreme during that time is not observed at any other site.
However, on average, core 77KL features positive values
of about +0.2% during H1, whereas sites 101KL and 114-3
show negative values around -0.2%. The transition from H1
into the B/A in core 77KL is characterized by decreasing
114-3 show average values of 0% and -0.2%, respectively,
with only low and insignificant amplitude variability. Overall, relative changes in d18Oivc-sw are regionally different,
and for the Bering Sea sites opposing trends with a maximum gradient of ~0.4% are observed. However, at each site,
a general covariation between d18Oivc-sw and TMg/Ca is
found. During 20–18 ka BP, cores 12KL and 77KL show
on average positive values (0% to +0.4%), whereas
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
Figure 4. Temperature gradients (ΔT) between SSTUk0 37 and TMg/Ca, and concentrations of CaCO3 and
biogenic opal for cores from the Bering Sea (SO201-2-77KL, -85KL, -101KL), the NW Pacific off
Kamchatka (SO201-2-12KL), and the southern Okhotsk Sea (LV29-114-3). Data are smoothed by
running five-point-averages (black lines). Dashed lines indicate ΔT = 0 C. Green vertical bars next to
the CaCO3 and opal axes mark respective concentrations in surface sediments in the direct vicinity of
the sediment core sites (stations SO201-2-11MUC, -76MUC, -83MUC, and -99MUC; Supplementary
Information). For comparison, records of d15Ndb from southern Okhotsk Sea core GGC27, southern
Bering Sea core HLY02-02-17JPC, and open subarctic Pacific core PC13 are shown (data from Brunelle
et al., 2007, 2010; cf. Figure 1). Pale orange and pale blue shadings represent the B/A and PB, and the H1
and YD, respectively.
d18Oivc-sw, while cores 12KL and 101KL show increasing
values. At Site 12KL, this increase is reflected by a positive
shift of about +1.4% occurring in less than 2 kyr (Figure 3B).
During the B/A, Site 77KL shows negative values of about 0.2% prevailing until the onset of the PB. In contrast, positive
values are recorded at sites 101KL (+0.3%) and 12KL (+1%),
where the heaviest deglacial d18Oivc-sw values are found. This
core shows a decrease in d18Oivc-sw during the B/A until a
minimum of -0.2% at the onset of the YD, and a subsequent
increase of 0.5% at 12.0 ka BP. Furthermore, core 12KL
recorded a pronounced maximum of +0.5% during the PB
(~11.5 ka BP), only found at this site, where it is subsequently
followed by a sharp decrease to about -0.5% until ~10 ka BP
(Figure 3B). Core 77KL also shows decreasing d18Oivc-sw
values since the onset of the PB until a minimum of -0.8%
is reached at 10–9 ka BP. During the Holocene, d18Oivc-sw
values are negative at all sites varying between -0.1% and 0.8%. Core 77KL from 9 to 6 ka BP shows a strong variability
with on average increasing d18Oivc-sw to about -0.1%. Even
lower values are subsequently preserved in this core until
~4 ka BP but affected by low temporal resolution.
5. Discussion
5.1. Deglacial Variability of Temperature and Salinity
[21] Our reconstructions show synchronous changes in
TMg/Ca and d18Oivc-sw during the last deglaciation. These
changes primarily reflect variations between “warm/salty”
and “cold/fresh” subsurface waters. Core 114-3, at least during the H1 to B/A transition, appears to show the opposite
relationship, but variability of TMg/Ca and d18Oivc-sw is low
for this core and not significantly above measurement uncertainty. Hence, subsurface conditions at Site 114-3 might
have remained fairly constant. Regional differences in the
deglacial d18Oivc-sw evolution are found, implying saltier
subsurface conditions during the B/A and PB in the NW Pacific (Site 12KL), while the Bering Sea cores show opposing
trends along the core transect. In contrast to sites 12KL and
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
during H1 [Gebhardt et al., 2008]. Hence, the northern
Emperor Seamounts might provide a different oceanographic setting when compared to the far NW Pacific and
the N Pacific marginal seas. Prior to ~15 ka BP, nondeterminable alkenone concentrations in our cores indicate
restricted alkenone production by coccolithophores, most
likely due to temperature limitation. Overall low marine productivity is supported by low biogenic opal concentrations.
At the same time, light diatom-bound nitrogen isotope ratios
(d15Ndb) in the subarctic Pacific realm [Brunelle et al., 2007,
2010] are indicative of a decrease in nitrate utilization in
response to the enhanced supply of preformed nutrients from
below and limited phytoplankton growth (Figure 4). Qualitative measurements of the sea-ice-related IP25 proxy imply
sea-ice influence at our sites during H1 and the YD, but
absent or at least considerably reduced influence during the
B/A and PB [Max et al., 2012]. Gebhardt et al. [2008]
argued for pronounced phases of sea-ice formation in the
subarctic N Pacific that induced vertical mixing during H1.
Evidence for increased vertical mixing and/or intensified
overturning during H1 and the YD comes from reduced
ventilation ages [Ohkushi et al., 2004; Sarnthein et al.,
2007; Sagawa and Ikehara, 2008; Okazaki et al., 2010],
and is supported by climate modeling studies [e.g., Okazaki
et al., 2010; Chikamoto et al., 2012; Menviel et al., 2012].
Moreover, benthic d18O excursions recorded in the northern
Bering Sea imply that the Bering Sea was a proximate
source of intermediate water during past stadial episodes
[Rella et al., 2012]. Increased ventilation and the potential
disappearance of the permanent halocline [Menviel et al.,
2012] as a result of higher salinity is consistent with the
reconstructions of Sarnthein et al. [2006] and Sagawa and
Ikehara [2008]. It is presumably also supported by our study
despite the regional differences in d18Oivc-sw, since results
for sites 77KL and 114-3 indicate the presence of saltier
subsurface waters (Figure 3B). The gradient in d18Oivc-sw along
our Bering Sea core transect rather reflects regional differences
in sea-ice growth and according brine rejection. We propose
that in the north, salty but 18O-depleted brines mixed with
subsurface waters explaining the lower d18Oivc-sw values with
respect to Site 77KL.
5.1.2. Bølling-Allerød
[24] Our records show maxima in TMg/Ca at the onset or
during the B/A and a subsequent cooling until the YD
(Figure 3A). Considering sea–ice influence, more positive
d18Oivc-sw values until ~14.0 ka BP imply enhanced sea-ice
melt or at least reduced sea-ice growth in the northern
Bering Sea, whereas negative and decreasing d18Oivc-sw
values were recorded at sites 77KL and 114-3 indicating
fresher subsurface conditions (Figure 3B). The recorded
evolution of TMg/Ca and SSTUk0 37 at our sites is congruent
with results from the NW Pacific [Sagawa and Ikehara,
2008], the NE Pacific [Kienast and McKay, 2001; Barron
et al., 2003; Gebhardt et al., 2008], the Bering Sea [Caissie
et al., 2010], and the Okhotsk Sea [Ternois et al., 2000;
Harada et al., 2006a; Seki et al., 2009]. The high TMg/Ca
of ~9 C at Site 12KL resemble conditions that today prevail
in the Kuroshio Extension area (e.g., at WOA station 32070)
[Locarnini et al., 2010] (cf. Figure 1), which, thus, argues
for the northward expansion of the N Pacific gyre filled with
Kuroshio waters as suggested by Gebhardt et al. [2008]. At
the same time, increased N Pacific ventilation ages [Adkins
101KL, sites 114-3 and 77KL indicate a freshening during
the B/A. However, all cores indicate a freshening since the
early Holocene. The deglacial SSTUk0 37 evolution is
different from TMg/Ca and characterized by warmings during
the B/A and PB, coolings during the H1 and YD, and a subsequent cooling during the early Holocene. Hence, changes
in SSTUk0 37 and N Atlantic climate occur quasi-synchronous
[Max et al., 2012]. The developing gradients between
surface and subsurface temperatures indicate changes in
thermal MLS, which were different on a regional scale.
[22] Strong amplitude variations in TMg/Ca and d18Oivc-sw
characterize Site 12KL, which might be explained oceanographically in the context of the major Bering Strait reopening. Today, the Alaskan Stream provides the NW Pacific and
the Bering Sea with relatively fresh surface waters. Some
authors speculated on reduced net inflow of Alaskan Stream
waters into the Bering Sea when some of the Aleutian passes
were closed due to lower glacial sea level [e.g., Katsuki and
Takahashi, 2005; Tanaka and Takahashi, 2005]. The Bering
Strait reopened between 12 and 11 ka BP [Keigwin et al.,
2006]. Accordingly, the inflow of waters from the Alaskan
Stream into the Bering Sea might have been reduced until
the PB allowing for relatively enhanced accumulation of
these water masses in the NW Pacific with respect to today.
During H1 and the YD, this should consequently have
resulted in more invigorated sea-ice formation and brine
rejection (providing 18O-depleted water) in the northern
Bering Sea, as well as in enhanced accumulation of cold/
fresh Alaskan Stream waters in the NW Pacific. In contrast,
during the B/A warm phase, sea-ice growth should have
been reduced (providing less 18O-depleted brines), and relatively warmer and saltier waters should have accumulated in
the NW Pacific. The weaker response at Site 12KL during
the PB is then explained by higher sea level and an open
Bering Strait, which resulted in stronger inflow of Alaskan
Stream waters into the Bering Sea.
5.1.1. Heinrich Stadial 1
[23] The above given considerations explain the presence
of cold/fresh subsurface waters at Site 12KL during H1
(Figure 3). Gebhardt et al. [2008] also related drops in N
Pacific salinity to North American river discharge and its
subsequent transport via the Alaskan Current and Alaskan
Stream rather than meltwater discharge from glacial ice
sheets. This is in agreement with the absence of a Beringian
Ice Sheet during the LGM [e.g., Brigham-Grette et al., 2001,
2003; Karhu et al., 2001]. Reduced inflow of Alaskan
Stream waters into the Bering Sea can also account for the
positive d18Oivc-sw at Site 77KL, implying relatively
increased subsurface salinity. At ODP Site 883D and at
Site MD01-2416 in the NW Pacific, SSTs were higher
than at our sites and characterized by three sharp increases
of ~4–6 C [Sarnthein et al., 2006; Gebhardt et al., 2008].
These SST pulses are accompanied by locally increased salinity and were explained by short-term incursions of
warm/salty Kuroshio waters. Warm SSTUk0 37 (20–24 C) of
Kuroshio waters are confirmed for the last deglaciation
[Sawada and Handa, 1998]. At Site GH02-1030, reconstructions from Globigerina bulloides show relatively saline
conditions prevailing since the LGM, with a maximum at
15.5 ka BP and a subsequent decrease in both salinity and
TMg/Ca [Sagawa and Ikehara, 2008]. NE Pacific core
MD02-2489 does also not show significantly higher SSTs
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
and Boyle, 1997; Ahagon et al., 2003; Ohkushi et al., 2004;
Sagawa and Ikehara, 2008; Okazaki et al., 2010] imply reduced formation of deep/intermediate water masses. The
presence of alkenones in all cores suggests that temperature
limitation of coccolithophores was overcome. ΔT values of
0–1 C indicate weak thermal MLS at sites 85KL and
101KL due to relatively stronger sea-ice influence and still
weak seasonal contrasts, which is supported by opal concentrations remaining low (<5 wt %; Figure 4). Sites 77KL and
114-3 also recorded only little higher opal concentrations.
However, TMg/Ca at these sites compare with those of sites
85KL and 101 KL, whereas SSTUk0 37 estimates are higher
by 2–3 C (Figure 3A). This indicates more pronounced
thermal MLS as a result of stronger seasonal contrasts. We
propose that continuing deglacial sea-level rise resulted in
enhanced net inflow of Alaskan Stream waters into the
Bering Sea providing Site 77KL with relatively fresher subsurface waters, thereby explaining the decrease in d18Oivc-sw
from H1 into the YD (Figure 3B). Relatively fresher subsurface conditions at sites 77KL and 114-3 are consistent with
surface freshening off Japan from 14.6 ka BP to 13.6 ka BP
[Sagawa and Ikehara, 2008]. Previous studies conducted
in the southern Bering and Okhotsk seas suggested
surface freshening due to invigorated sea-ice melt from
the observation of negative shifts in planktonic d18O
and increasing abundances of the diatom species Paralia
sulcata [Gorbarenko et al., 2004, 2005; Seki et al.,
2004a]. In contrast, opal concentrations at Site 12KL of
up to ~12 wt % are comparable with Holocene values
and suggest significantly enhanced marine productivity.
At the same time, heavy d15Ndb were recorded in the
southern Bering Sea and NW Pacific [Brunelle et al.,
2010], implying enhanced nitrate utilization or even denitrification of nitrate as a result of stronger upper-ocean
stratification (Figure 4). The negative ΔT values in core
12KL during the B/A could be related to a significantly
warmer and deepened mixed layer during a prolonged
summer season and reduced sea-ice influence. Consequently, for Site 12KL, we consider seasonal bias in
SSTUk0 37 signal formation, and that during the B/A,
SSTUk0 37 more likely represent early spring conditions.
Coccolithophorid blooms that in the modern subarctic N
Pacific occur already during spring are known [Takahashi
et al., 2002], and for the Okhotsk Sea seasonal bias has
been suggested before to explain similarly high LGM
and Holocene SSTUk0 37 values [Seki et al., 2004b].
5.1.3. The Younger Dryas and Preboreal
[25] The TMg/Ca and d18Oivc-sw evolution observed at sites
114-3, 12KL, and 77KL during the YD and PB (Figure 3)
broadly matches that found in the NW Pacific [Sarnthein
et al., 2004, 2006; Gebhardt et al., 2008; Sagawa and
Ikehara, 2008], although our records show lower d18Oivc-sw
values. Lowered SSTUk0 37 during the YD are in agreement
with the suggested southeastward expansion of the W Pacific
Subarctic Gyre and the Bering Sea Gyre [Kienast and McKay,
2001], as well as with modeling results predicting a strengthened and eastward expanded Aleutian Low, which initiated
better ventilation [Mikolajewicz et al., 1997]. This is also in
agreement with reduced N Pacific ventilation ages [e.g.,
Okazaki et al., 2010]. At the same time, similar TMg/Ca and
SSTUk0 37 indicate weak but regionally different thermal MLS
at our sites as a result of less developed seasonal contrasts.
At sites 12KL, 77KL, and 114-3, ΔT values of ~2 C imply relatively stronger thermal MLS than at Site 85KL (ΔT = 0 C),
which recorded H1-like conditions (Figures 3 and 4). During
the PB, TMg/Ca and related d18Oivc-sw values imply further subsurface cooling and freshening (77KL) and/or conditions still
comparable to the YD (114-3). Concurrent increases in
SSTUk0 37 argue for more pronounced thermal MLS, which
is characterized by a deepening of the thermocline, extra
cooling of sub-thermocline waters in hand with the formation of the dichothermal layer. We propose that at least
our Bering Sea sites were now subject to increased inflow
of relatively warmer Pacific surface waters through the
Aleutian passes as a result of the reopening of the Bering
Strait, and that this process fostered seasonal MLS by
reducing the sea-ice growth.
5.1.4. The Holocene
[26] Most notably for all our cores are the differently
developing TMg/Ca and SSTUk0 37 since the PB (Figure 3A),
best explained by insolation changes affecting the surface
ocean warming, reduced sea-ice, and increasing seasonal
contrasts controlling the shape and position of the thermocline. Holocene TMg/Ca remain either equal or stay lower
than during H1 (Figure 3A). Our TMg/Ca, although being
1–2 C warmer, well reflect modern conditions, which is
consistent with the NW Pacific studies [Sarnthein et al.,
2004, 2006; Gebhardt et al., 2008; Sagawa and Ikehara,
2008]. Overall, d18Oivc-sw values remain negative at all sites
indicating relatively fresh subsurface conditions and almost
stable seasonal contrasts. Sarnthein et al. [2004] reported
on the long-term but stepwise decrease in subsurface salinity
at Detroit Seamount during the early to middle Holocene,
speculating that the modern salinity-driven stratification developed only since the Holocene. We can not confirm this
stepwise salinity decline. However, our results point to a
deglacial evolution of subsurface water mass characteristics
developing into relatively fresh conditions along with enhanced thermal MLS (strong seasonal contrasts) since the
early Holocene. Dominant control on the Holocene subarctic
NW Pacific SSTUk0 37 evolution is attributed to Northern
Hemisphere summer insolation [Okumura et al., 2009; Hu
et al., 2010; Max et al., 2012]. In consequence, stronger seasonal contrasts developed as the result of a prolonged summer season and reduced sea-ice influence. Stronger winter
mixing would have lead to the formation of the dichothermal
layer. In this case, our high ΔT values of ~5–6 C would be
explained by higher SSTUk0 37 due to stronger insolation,
and by Nps recording gradually cooler subsurface
temperatures due to the presence of the dichothermal layer.
Accordingly, thermal MLS was consolidated and the
seasonal halocline developed from sea-ice melt during
summer. The largest difference between TMg/Ca and SSTUk0 37
at ~10 ka BP coincides with the early Holocene thermal
maximum, which in Alaska and northwest Canada
occurred between 11 and 9 ka BP [Kaufman et al.,
2004]. Recent model experiments indicate that besides
insolation the reopening of the Bering Strait might have
further affected this SSTUk0 37 maximum [Okumura
et al., 2009; Hu et al., 2010]. Additionally, enhanced
precipitation from the Westerlies and strengthened advection of cold/fresh waters from the Alaskan Stream might
have contributed to the N Pacific cooling and freshening
[Sarnthein et al., 2004].
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RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
5.2. Cause of Deglacial Stratification Changes and
Impact on Atmospheric CO2
[27] The deglacial SSTUk0 37 evolution recorded in the subarctic N Pacific realm appears to be related to and in-phase
with the thermal evolution recorded in the N Atlantic realm,
suggesting an atmospheric coupling [Ternois et al., 2000;
Kienast and McKay, 2001; Barron et al., 2003; Harada
et al., 2006a; Seki et al., 2009; Caissie et al., 2010; Max
et al., 2012]. Several modeling studies investigated the
sensitivity of the PMOC in response to perturbations of the
AMOC on millennial time scales during the last deglaciation. All models predicted an enhanced PMOC at times of
a weakened AMOC, whose strength was modulated by
freshwater input into the Atlantic or by freshwater extraction
from the Pacific [e.g., Saenko et al., 2004]. In the models,
the weakened AMOC during H1 and the YD [e.g.,
McManus et al., 2004] results in the southward shift of the Intertropical Convergence Zone and in weakened Indian and
Asian summer monsoons [Zhang and Delworth, 2005], which
today maintain the salinity gradient between the Atlantic and
Pacific [e.g., Emile-Geay et al., 2003]. Consequently, the antiphase models [e.g., Saenko et al., 2004] predicted a warming
in the N Pacific due to the weakened AMOC, whereas the
in-phase models [e.g., Mikolajewicz et al., 1997] suggested
that atmospheric cooling by a strengthened Aleutian Low
resulted in a PMOC. Interestingly, modeling results of
Chikamoto et al. [2012] showed a cooling in the N Atlantic
and in the NW Pacific, but ambiguous results for the NE Pacific. This was explained by a salinity increase in the N Pacific
due to the weakened east Asian monsoon and changes in the
oceanic transport of heat and salt. Both in-phase and antiphase
models are supported by low ventilation ages in the N Pacific
during H1 and the YD. Overall, our results support the studies
of Chikamoto et al. [2012] and of Menviel et al. [2012], who
discussed the effects of the removed halocline in the subarctic
N Pacific. However, we cannot confirm strongly enhanced
opal production, which was also predicted for H1 as a result
of the established PMOC and the removal of the halocline
[Menviel et al., 2012]. We rather attribute the low alkenone
and opal concentrations observed at our sites to limitation of
marine productivity by temperature and sea-ice. The notion
of the removed halocline during H1 is supported by the
high d18Oivc-sw values recorded in NW Pacific cores
MD01-2416 [Sarnthein et al., 2006; Gebhardt et al.,
2008] and GH02-1030 [Sagawa and Ikehara, 2008].
During H1, our reconstructions indicate cold or decreasing surface and subsurface temperatures (atmospheric
cooling), while our results for d18Oivc-sw indicate a subsurface freshening at NW Pacific Site 12KL (strong Alaskan
Stream). At the same time, saltier-than-today subsurface
conditions are recorded at Bering Sea Site 77KL (weaker
halocline), and increasing brine influence toward the north is
assumed. The TMg/Ca and d18Oivc-sw recorded at our sites for
the B/A and PB is explained by stronger seasonal MLS, which
might be attributed to the reduced PMOC. Subsequent to
the PB, the subarctic NW Pacific is subject to enhanced
thermal MLS, which for the Bering Sea sites is possibly
related to the opening of Bering Strait resulting in stronger
inflow of N Pacific surface waters. Moreover, increased
surface freshening during the Holocene potentially resulted
from enhanced sea-ice melting during summer and enhanced
precipitation.
[28] The deep N Pacific is suggested to have stored a
greater portion of respired CO2 during glacials than during
interglacials [Jaccard et al., 2009]. During the last deglaciation, when the AMOC was strong, the respired carbon pool
was removed from the deep sea, which is considered to have
resulted in a deepening of the lysocline and an increase in
atmospheric CO2 [e.g., Galbraith et al., 2007]. Our records
confirm increased CaCO3 preservation and higher marine
productivity during the B/A. Conversely, Okazaki et al.
[2010] argued for enhanced N Pacific deep water formation
during H1 (i.e., better ventilation when the AMOC was
weak), and Rella et al. [2012] suggested that the Bering
Sea was a proximate source for NPIW during that time. This
is in agreement with the notion of a removed halocline,
which would result in a rise in atmospheric CO2 of 5 ppm
[Chikamoto et al., 2012; Menviel et al., 2012]. Another
aspect involves the efficiency of the biological pump. As recently summarized for the Pliocene, the subarctic N Pacific
might have been responsible for higher atmospheric CO2
in case of deep water formation resulting in a reduced ratio
of regenerated-to-preformed nutrients in the ocean, thereby
lowering the efficiency of the biological pump [Studer et
al., 2012]. The d15Ndb data of Brunelle et al. [2007, 2010]
indeed argue for reduced nitrate utilization in the NW Pacific
and its marginal seas during H1 and the YD (Figure 4).
Since we reconstructed changes in thermal MLS, we can
neither argue for nor against a change in the position and/
or the strength of the permanent halocline. Nevertheless,
our d18Oivc-sw data show higher subsurface salinity in the
southern Bering Sea (Site 77KL) during H1, but during the
B/A in the NW Pacific (Site 12KL). This can at least be
taken as supportive evidence for regional changes in subarctic N Pacific permanent halocline stability, which might
therefore have contributed to the deglacial rise in atmospheric CO2.
6. Conclusions
[29] We found regionally different changes in TMg/Ca and
d18Oivc-sw in the subarctic NW Pacific as well as differences
in alkenone- and Mg/Ca-based SST reconstructions. Our
results indicate deglacial oceanographic changes in the
mixed layer, which are most likely related to variations in
Northern Hemisphere summer insolation and the strength
of atmospheric pressure systems. Both factors control the
extent of seasonal contrasts by driving changes in sea-ice
formation, thereby altering seasonal MLS. From our results,
we propose that seasonal contrasts and, hence, thermal MLS,
although being regionally different, were reduced during H1
and the YD, but strong during the B/A. Our sites were characterized by low marine productivity, cold subsurface
waters, and weak thermal MLS during the H1 and YD cold
phases. This is attributed to an established PMOC and atmospheric cooling resulting in limited phytoplankton growth
and enhanced sea-ice formation. Additional influence by
increased advection of Alaskan Stream waters accumulating
in the NW Pacific due to a closed Bering Strait is proposed.
In contrast, warmer subsurface waters and regionally different temperature gradients were recorded during the B/A. At
the same time, we found relatively increased d18Oivc-sw in
the NW Pacific (Site 12KL) and northern Bering Sea (Site
101KL), but reduced d18Oivc-sw in the southern Okhotsk
101
RIETHDORF ET AL.: DEGLACIAL SUBSURFACE TEMPERATURE AND SALINITY IN THE NW PACIFIC
and Bering seas (sites 114-3 and 77KL). This is explained
by enhanced thermal MLS, accumulation of Alaskan Stream
waters in the NW Pacific (at Site 12KL), as well as a stronger impact of brines on the d18Oivc-sw signal in the northern
Bering Sea. A decrease in TMg/Ca and d18Oivc-sw during the
early Holocene along with high temperature gradients
implies the only recent establishment of modern MLS in
the subarctic NW Pacific and argues for subsurface oceanographic changes, which might be related to the reopening of
the Bering Strait at 12–11 ka BP.
[30] Acknowledgements. This study was funded by the German
Federal Ministry of Education and Research (BMBF; grant 03G0672A
and B) and resulted from the joint German-Russian research project
“KALMAR.” Master and crew of R/V Sonne cruise SO201 Leg 2 are
gratefully acknowledged for their professional support. E. Maier conducted
additional opal measurements for core 77KL. We thank N. Gehre,
L. Haxhiaj, B. Domeyer, and D. Poggemann for laboratory assistance and
technical support, N. Khelifi for discussions, as well as two anonymous
reviewers and the Editor for considerably improving the manuscript. The
data are available via the PANGAEA Data Publisher for Earth & Environmental Science (http://doi.pangaea.de/10.1594/PANGAEA.786258).
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